onset of fault reactivation in the eastern cordillera of ...cmontes/camilomontes...palaeogene time...

Onset of fault reactivation in the Eastern Cordillera of Colombiaand proximal Llanos Basin; response to Caribbean–South

American convergence in early Palaeogene time

GERMAN BAYONA1*, AGUSTIN CARDONA1,2, CARLOS JARAMILLO1,3,

ANDRES MORA4, CAMILO MONTES1,5, VICTOR CABALLERO4,

HERNANDO MAHECHA1, FELIPE LAMUS1, OMAR MONTENEGRO1,

GIOVANNY JIMENEZ1, ANDRES MESA6 & VICTOR VALENCIA7

1Corporacion Geologica ARES, Calle 44A, No. 53-96, Bogota, Colombia2Universidad Nacional de Colombia, Medellın, Colombia

3Smithsonian Tropical Research Institute,Ciudad de Panama, Panama4Instituto Colombiano del Petroleo,Bucaramanga, Colombia

5Universidad de los Andes, Bogota, Colombia6Hocol S.A., Bogota, Colombia

7School of Earth and Environmental Sciences, Washington State University,

Pullman, WA 99164-2812 85721, USA

*Corresponding author (e-mail: [email protected])

Abstract: The inversion of Mesozoic extensional structures in the Northern Andes has controlledthe location of syn-orogenic successions and the dispersal of detritus since latest Maastrichtiantime. Our results are supported by detailed geological mapping, integrated provenance (petrogra-phy, heavy minerals, geochronology) analysis and chronostratigraphical correlation (palynologicaland geochronology data) of 13 areas with Palaeogene strata across the central segment of theEastern Cordillera. Spatial and temporal variation of sedimentation rates and provenance dataindicate that mechanisms driving the location of marginal and intraplate uplifts and tectonic sub-sidence vary among syn-orogenic depocentres. In the late Maastrichtian–mid-Palaeocene time,crustal tilting of the Central Cordillera favoured reverse reactivation of the western border ofthe former extensional Cretaceous basin. The hanging wall of the reactivated fault separatedtwo depocentres: a western depocentre (in the Magdalena Valley) and an eastern depocentre (pre-sently along the axial zone of the Eastern Cordillera, Llanos foothills and Llanos Basin). In latePalaeocene–early Eocene time, as eastern subduction of the Caribbean Plate and intraplate mag-matics advanced eastwards, reactivation of older structures migrated eastwards up to the LlanosBasin and disrupted the eastern depocentre. In early Eocene time, these three depocentres were sep-arated by two low-amplitude uplifts that exposed dominantly Cretaceous sedimentary cover. Syn-orogenic detrital sediments supplied from the eastwards-tilted Central Cordillera reached areas ofthe axial domain of the Eastern Cordillera, whereas unstable metamorphic and sedimentaryfragments recorded in the easternmost depocentre were supplied by basement-cored uplifts withCretaceous and Palaeozoic sedimentary cover reported in the southern Llanos Basin. This tec-tonic configuration of low-amplitude uplifts separating intraplate syn-orogenic depocentres andintraplate magmatic activity in Palaeocene time was primary controlled by subduction of theCaribbean Plate.

Supplementary material: Appendix 1 presents detailed descriptions of analytical methods used inthis manuscript. Appendixes 2 to 4 include raw data of sandstone petrography, heavy minerals andU–Pb detrital zircon geochronology, respectively. All this material is available at http://www.geolsoc.org.uk/18597.

Reactivation of pre-existing structures occurs asone response to changes in tectonic regime. Posi-tive inversion structures (Hayward & Graham 1989;Coward 1994; Sibson 1995; Bayona & Lawton

2003) involve the reactivation of a former normalfault; in the Northern Andes, major uplifts haveoccurred along inversion structures (e.g. Cooperet al. 1995; Colleta et al. 1997; Cortes et al. 2006;

From: Nemcok, M., Mora, A. R. & Cosgrove, J. W. (eds) Thick-Skin-Dominated Orogens: From InitialInversion to Full Accretion. Geological Society, London, Special Publications, 377,http://dx.doi.org/10.1144/SP377.5 # The Geological Society of London 2013. Publishing disclaimer:www.geolsoc.org.uk/pub_ethics

10.1144/SP377.5 Geological Society, London, Special Publications published online March 8, 2013 as doi:

http://www.geolsoc.org.uk/18597



Mora et al. 2008a). Several factors may controlwhether an older structure may be reactivated: (a)fault and basin geometries (Butler 1989; Bucha-nan & McClay 1991, 1992; Bayona & Thomas,2003); (b) potential of fault reactivation (e.g. porepressure, friction, cohesion and fault gouge mate-rial: Sibson 1985, 1995; Etheridge 1986); and (c)reorganization of stresses along the collided margin(Kluth & Coney 1981; Ziegler et al. 1998; Marshaket al. 2000).

Inversion structures fragment former basin geo-metry and drive the evolution of compartmentedsyn-orogenic basins. As fault reactivation devel-ops, tectonic fragmentation of the drainage system,depositional patterns and climate conditions beginto differ between localized depocentres (Bayona& Thomas 2003; Sobel & Strecker 2003; Sobelet al. 2003; Hilley & Strecker 2005; Strecker et al.2007, 2009, 2011). Therefore, depositional envi-ronments, stratigraphical thicknesses and prove-nance signatures begin to vary among depocentres(Bayona & Lawton 2003; Hain et al. 2011). As aresult, the Magdalena Valley Basin and the Llanosforeland basin of Colombia are two basins withdrainage, climate and depositional systems thatwere totally disrupted by the rapid vertical upliftof the Eastern Cordillera of Colombia that occur-red since late Miocene–Pliocene time (Fig. 1a)(see Mora et al. 2008b for review of the youngestevents of Eastern Cordillera uplift). However,upper Cretaceous rocks of the Magdalena Valley,Eastern Cordillera and Llanos Basin accumulatedwithin the same basin and represent a single depo-sitional profile of a shallow platform in a passive-margin setting (Cooper et al. 1995; Villamil 1999;Cediel et al. 2003).

Conversely, modelling of undisrupted basins,such as foreland basins (DeCelles & Giles 1996),has allowed predictions of patterns of deforma-tion (i.e. the advance of orogenic front and flexuralsubsidence), sediment routing (provenance signa-ture) and accumulation patterns (stratal geometryand three-dimensional lithofacies distribution) tobe made (Beaumont 1981; Jordan 1981; Helleret al. 1988; Flemings & Jordan 1989; Crampton &Allen 1995; Burbank et al. 1996; DeCelles &Giles 1996; Sinclair 1997; DeCelles & Horton2003). In orogens with multiphase deformation,such as the Northern Andes (Bayona et al. 2008),comprehensive basin stratigraphy and provenanceanalyses become the key to document the timingand sense of fault movement along reactivatedstructures.

The timing of when a basin becomes a positivefeature is relevant for the determination of the ces-sation of burial processes of source rocks and theformation of early structures in a petroleum sys-tem. Definition of the onset of positive inversion

in the Northern Andes has been controversial, andthe timings range from Maastrichtian–Palaeocene(Sarmiento-Rojas 2001; Montes et al. 2005; Corteset al. 2006; Bayona et al. 2008; Moretti et al. 2010;Parra et al. 2012), to middle–late Eocene (Toro et al.2004; Gomez et al. 2005a; Parra et al. 2009a, b;Horton et al. 2010b; Saylor et al. 2011, 2012), lateEocene (Mora et al. 2010) and late Oligocene–Miocene (Cooper et al. 1995; Villamil 1999). Inorder to address this problem of the timing offault reactivation in a multiphase orogeny, like theEastern Cordillera, we integrated the analysis ofdepositional environments, thickness variationsand provenance signatures of Maastrichtian–lowerEocene strata that are presently compartmentalizedby several structures within the double-vergingEastern Cordillera of Colombia (Fig. 1). Our studyspecifically examines the Maastrichtian–early Pal-aeogene time interval to document possible intra-plate deformation events coeval with the obliqueconvergence of the Caribbean Plate or the collisionof oceanic arcs with the South American Plate (seethe review in Bayona et al. 2012). The approachused in this study may be applied in multiphaseorogens where low-temperature thermochronome-try does not reveal early phases of shortening butabrupt changes in provenance signatures and sedi-mentation rates record those early phases ofdeformation.

Regional setting

Three major mountain belts are the result of thecomplex interaction of the Nazca, Caribbean andSouth America plates since the Late Cretaceous:Western Cordillera (WC), Central Cordillera (CC)and the double-verging Eastern Cordillera (EC)(Fig. 1a, b). The Eastern Cordillera is interpretedas a wide Cretaceous extensional basin formedduring at least two stretching events (Sarmiento-Rojas et al. 2006) and tectonically inverted duringthe Cenozoic (Colleta et al. 1990; Dengo & Covey1993; Cooper et al. 1995; Mora et al. 2006). Thegeometry of the double-verging Eastern Cordillerachanges northwards from a narrow belt in thesouthern domain that broadens in its central seg-ment where the Cordillera is bounded by east- andwest-verging thrust belts (Fig. 1b). Further to thenorth, the Eastern Cordillera bifurcates into theSantander Massif and the Merida Andes (Fig. 1a).

The areas selected for this study are located inthe three structural domains of the central segmentof the Eastern Cordillera (Fig. 1c). Its westerndomain includes a west-verging thrust and foldbelt that places Cretaceous and Palaeogene strata incontact with Cenozoic strata and buried structuresof the southern Middle Magdalena Valley Basin. In

G. BAYONA ET AL.10.1144/SP377.5

Geological Society, London, Special Publications published online March 8, 2013 as doi:

the frontal splay of the west-verging system is thenorth–south-trending Guaduas Syncline (Corteset al. 2006). The axial domain consists of tight anti-clines and broad synclines related to double-vergingfaults (McLaughlin & Arce 1975; Fajardo-Pena1998; Toro et al. 2004; Bayona et al. 2008). This

domain broadens northwards of Bogota City, andis divided into western axial and eastern axialdomains by the east-verging Pesca–Macheta FaultSystem and the Boyaca–Soapaga Fault System(Fig. 1c). The eastern domain includes east-vergingfaults placing Cretaceous and Cenozoic strata over

Fig. 1. (a) Location of the three Cordilleras in the Northern Andes and Romeral Fault System (RFS). (b) Regionalcross-section from the Northern Andes (modified from Bayona et al. 2012). (c) Geological map showing majorstructures of the Eastern Cordillera and the location of selected studied areas. For reference, the studied areas weregrouped into four structural domains: western; western axial; eastern axial; and eastern domains.

ONSET OF FAULT REACTIVATION10.1144/SP377.5


Cenozoic strata of the Llanos foreland basin. Base-ment rocks of the Quetame Massif are exposed inthe southern segment of the eastern domain (Fig.1c). The NE-trending Medina Syncline is part ofthe frontal splay of the east-verging fault system(Parra et al. 2009a). Reverse structures in thesethree domains of the Eastern Cordillera have beeninterpreted as an inversion of former normal faults(Cooper et al. 1995; Branquet et al. 2002; Restrepo-Pace et al. 2004; Cortes et al. 2006; Kammer &Sanchez 2006; Mora et al. 2006, 2008a).

Interpretations of basin geometry for the earlyPalaeogene vary from a single and continuous fore-land basin (Cooper et al. 1995; Gomez et al. 2005a;Parra et al. 2009b; Horton et al. 2010b; Sayloret al. 2011, 2012), a continuous negative flexuralbasin with an absence of bounding thrusts (Pindellet al. 2005) to a foreland basin disrupted by uplifts(Bayona et al. 2008). The location of intraplateuplifts has been proposed in the western border ofthe basin (Bayona et al. 2003; Restrepo-Pace et al.2004; Cortes et al. 2006; Moretti et al. 2010),along the axial zone of the basin (Fabre 1981,1987; Sarmiento-Rojas 2001; Pardo 2004) or atboth borders of the basin (Fajardo-Pena 1998; Villa-mil 1999). Geodynamic modelling of Palaeogenestrata encompassing the axial zone of the EasternCordillera to the Llanos Basin indicates an east-wards migration of tectonic loads within therestored Eastern Cordillera to explain the Maas-trichtian–Palaeogene geometry of the syn-orogenicbasin (Bayona et al. 2008).

Early Palaeogene (45–65 Ma) magmatic activityhas been reported along the northern Andean conti-nental margin (Central Cordillera–Santa MartaMassif in Fig. 1a, b) (Aspden et al. 1987; Cardonaet al. 2011), and volcanic detrital zircons havebeen recognized in Palaeocene–lower Eocene sand-stones as far as 400 km from the magmatic arc(Saylor et al. 2011, 2012; Bayona et al. 2012).Marginal and intraplate magmatism during earlyPalaeogene time is related to the convergence andshallow subduction of the Caribbean crust beneaththe South America Plate, and the abrupt termi-nation of volcanism in middle Eocene time to achange to a transpressive non-magmatic margin(Bayona et al. 2012).

Maastrichtian–lower Eocene stratigraphy

and depositional environments

In this section we summarize lithological associ-ations and depositional environment interpretationsproposed for Maastrichtian–lower Eocene unitsin the three structural domains of the centralsegment of the Eastern Cordillera, as these unitsrecord the regional shift from marine to continental

depositional environments as a response to theaccretion of oceanic terranes west of the RomeralFault System (Fig. 1) (e.g. Etayo-Serna et al.1983; McCourt et al. 1984).

Maastrichtian–lower Eocene stratigraphicalunits of the western domain have been definedin the western flank of the Guaduas Syncline (Fig.1c), and include, from base to top, the Umir, Cimar-rona, Seca and Hoyon formations (Fig. 2). Thelower two units change laterally and vertically fromcalcareous dark-grey mudstones with abundant fos-sil content (bivalves, foraminifera), to cross-beddedconglomerates and fine-grained sandstones (Gomez& Pedraza 1994); conglomerate beds of the Cimar-rona Formation change eastwards to limestonebeds in the eastern flank (Gomez et al. 2003). Theoverlying Seca Formation includes dark-colouredmudstones with thin coal interbeds at the base,changing upsection to varicoloured massive mud-stones with interbeds of fine- to medium-grainedsandstones towards the top (De Porta 1966). TheHoyon Formation comprises interbeds of mud-stones, sandstones and conglomerates that range insize from metre- to decametre-scale, and is dividedinto three informal units. The lower Hoyon unit is aconglomerate interval with upwards-coarseningsuccessions, and has transitional contacts with sand-stones and mudstones of the underlying Seca For-mation and with mudstones of the overlyingmiddle Hoyon unit; the latter units comprise domi-nantly massive red mudstones. In contrast, theupper Hoyon unit is a conglomerate interval rest-ing unconformably upon the middle Hoyon unit.Both in the lower and the upper Hoyon units, hori-zontal bedding dominates over cross-bedding asinternal structures, and an east–NE direction oftransport is documented by palaeocurrent measure-ments (Gomez et al. 2003). Gomez et al. (2003)interpreted the Maastrichtian–lower Eocene suc-cession as a regression from a shallow-marineenvironment with fan deltas (the Umir and Cimar-rona formations), changing to coastal and allu-vial plains (the Seca Formation), and ending upwith fluvial and alluvial fan deposits (the HoyonFormation).

In the axial domain, Maastrichtian–lower Palae-ocene strata include the Guaduas Formation, over-lain by Palaeocene–lower Eocene Cacho andBogota formations to the west of the Pesca–Macheta Fault System, or by the Lower and UpperSocha formations to the east of the Pesca–Macheta Fault (Fig. 2). The Guaduas Formationrests conformably over fine- to medium-grainedsandstones of the Guadalupe Formation, with cross-beds and abundant ichnofossils (Perez & Salazar1978; Fabre 1981). The lower half of the GuaduasFormation includes fine-grained strata with abun-dant coal seams, whereas the upper half includes



laminated mudstones, and massive light-colouredmudstone and sandstone interbeds (Sarmiento1992). The Cacho–Lower Socha formations arecross-bedded, fine- to coarse-grained upwards-fining

quartzarenites, locally conglomeratic; these unitsrest disconformably upon light-coloured mudstonesof the Guaduas Formation (Pardo 2004; Bayonaet al. 2010). In the western axial domain, the

Fig. 2. (a) Lithostratigraphical correlation of the Maastrichtian–Lower Eocene strata in the four structural domains,showing lateral variation of the thickness and lithology of stratigraphical units (the names of the lithological units areshown to the left of each structural domain). Datum, lower–middle Eocene unconformity. (b) Chronostratigraphicalcorrelation of upper Campanian–middle Eocene strata based on palynological and geochronological data. See Figure 1for the location of sections, and Table 1 for a summary of thickness and chronostratigraphical data per area.



Bogota Formation overlies conformably the CachoFormation, and consists dominantly of massive mul-ticoloured sandy mudstone with isolated upwards-fining and -coarsening massive to cross-beddedsandstones (Bayona et al. 2010). In the easternaxial domain, the Upper Socha Formation over-lies conformably the Lower Socha Formation andit changes upsection from bioturbated dark-greyorganic-rich claystone and mudstone with thincoal seams to massive light-coloured sandy mud-stone with isolated upwards-fining and -coarseningsandstones with cross beds, wavy lamination andripple cross-lamination (Pardo 2004). Coarse-grained to conglomeratic sandstones of the Rega-dera–Picacho formations disconformably overliethe Bogota–Upper Socha formations. Sarmiento(1992) interpreted the regression from shallow-marine to alluvial plain conditions for the GuaduasFormation, whereas Bayona et al. (2010), Pardo(2004) and Saylor et al. (2011) interpreted over-lying units as deposition in fluvial systems varyingfrom braided, to meandering to anastomosing.

In the eastern domain, as well as in the proxi-mal Llanos Basin, the Guaduas Formation is verythin or absent, and the Barco and Cuervos for-mations rest unconformably on upper Cretaceousunits (Cooper et al. 1995; Bayona et al. 2008) (Fig.2). The succession described below correspondsto the western flank of the Medina Syncline. Maas-trichtian strata consist of a coarsening-upwards suc-cession from medium-grained to conglomeraticsandstones with cross-beds and reworked bivalves,oysters and corals of the Guadalupe Formation(Guerrero & Sarmiento 1996). The Barco Forma-tion comprises fine- to medium-grained quartzosesandstones and is overlain by interbeds of upwards-fining sandstone and mudstone with bidirectionalcross-bedded and bioturbated heterolithic lami-nated sandstone of the lower Cuervos Formation.The middle and upper Cuervos Formation resem-bles the lithological association of the UpperSocha Formation. Guerrero & Sarmiento (1996)interpreted Maastrichtian strata as shallow-marinedeposits, whereas Cazier et al. (1995), Reyes(1996) and Bayona et al. (2008) interpreted theBarco–Cuervos succession as coastal plains with atidal influence. The Mirador Formation is a coarse-grained sandstone unit of lower–middle Eoceneage (Bayona et al. 2008; Parra et al. 2009a) thatdisconformably overlies the Cuervos Formation.

Selection of study areas and methods

In order to determine spatial and temporal changesin basin geometry for a given time interval, it iscrucial to consider the chronostratigraphical cor-relation of lithological units described earlier

at different structural domains. The survey of Maas-trichtian–lower Eocene strata included three areasin the Guaduas Syncline (San Juan de Rıo SecoEast and West, and Guaduero), six areas in the west-ern axial zone (Usme, Guatavita, Macheta, Checua,Tunja and Paipa), five areas in the eastern axial zone(Umbita, San Antonio, Rondon, Pesca and Paz deRıo) and two areas in the eastern domain (Cusianaand Medina) (Fig. 1b).

The thickness of the Maastrichtian–lowerEocene lithostratigraphical units described previ-ously was measured in synclinal structures andwithin areas with minimum structural complex-ity. Lithostratigraphical units were mapped in bothflanks of each syncline structure, and the lateralvariation of thickness was calculated using localstructural cross-sections. Samples for palynologi-cal and provenance analyses (petrography, heavymineral and detrital U–Pb zircon geochronology)were collected for each measured stratigraphicalsection.

Palynology and detrital volcanic zircons sup-ply the age control for Maastrichtian–lower Eo-cene units. Descriptions of analytical methods andraw data for palynological, sandstone petrography,heavy minerals and geochronological analysesare presented in the Supplementary Material. Paly-nology and detrital volcanic zircons supply theage control for Maastrichtian-lower Eocene units.The results of 381 palynological samples are sum-marized in Table 1. Calculation of maximum ageof deposition based on U–Pb ages in volcaniczircons follows the results presented in Bayonaet al. (2012).

Sandstone petrography (123 samples), heavymineral (46 samples) and U–Pb detrital zircongeochronology (28 samples) analyses were carriedout in late Campanian–lower Eocene sandstonesin order to compare composition, heavy mineralassemblage and detrital zircon age populationsbetween structural domains. We integrate theresults of seven U–Pb detrital zircon geochronologysamples for Maastrichtian–Lower Eocene stratain the Floresta Basin (or Paz de Rıo area of theeastern axial domain: Saylor et al. 2011, 2012),one sample of lower Eocene strata in the Machetaarea (Horton et al. 2010b) and two samples of thenorthern Guaduas Syncline (Caballero et al. 2013)in order to compare detrital zircon age populationsfrom the western, axial and eastern domains fromCampanian to early Eocene times.

Heavy minerals were grouped into ultrastable(zircon, rutile, turmaline), stable (apatite, garnet),unstable (amphibole, epidote, pyroxene, chlorite,talc and others) and muscovite. Zircon crystalli-zation ages are divided into three major prov-inces: 65–360 Ma from the Central Cordillera;360–1300 Ma from basement rocks of the Eastern



Cordillera and Santander Massif; and zircons sup-plied from basement rocks in the Guyana Cratonhave Neoproterozoic ages (600–650 Ma) and popu-lations older than 1300 Ma (see Horton et al. 2010a,b for a detailed discussion of zircon crystallizationages in the Northern Andes). The Cretaceous sed-imentary cover of the Eastern Cordillera includesa mixture of the last two provinces (see detailslater).

Thickness variation and age control of

Maastrichtian–lower Eocene strata

This section presents the most significant trendsof stratigraphical thickness variations for studiedstratigraphical units within structural domains.Figure 2a illustrates the measured stratigraphicalthickness for each unit in each area and the positionof samples with volcanic detrital zircons. Palynolo-gical results (summarized in Table 1) and calcu-lations of maximum depositional age were used toconstruct the chronostratigraphical correlation pre-sented in Figure 2b.

In the western domain, the stratigraphical thick-ness of the Seca Formation increases eastwardsand northwards (Table 1; Fig. 2a). In the area ofGuaduero, the thickness ranges from 490 m tomore than 700 m, whereas in the San Juan de RioSeco area the thickness ranges from 400 to 470 m.The lower and middle intervals of the Hoyon For-mation are completely recorded in the eastern flankof the San Juan de Rio Seco area, and are partlyeroded by the unconformity in the western flankof the San Juan de Rio Seco area and completelyeroded in the Guaduero area (Fig. 2a). For theSeca Formation, a Maastrichtian palynological ageat the base and detrital volcanic zircon ages at thetop (61.9 + 2 Ma) constrain a Maastrichtian–lower Palaeocene age (Table 1). The lower Hoyonunit yielded only two zircons with a Palaeoceneage, 62 Ma being the youngest zircon, whereas themaximum age of deposition for the middle Hoyonunit is 56.3 + 1.6 Ma. The upper Hoyon yieldedmiddle Eocene–lower Oligocene palynologicalages, and only one zircon yielded a 55 Ma age.

In the western axial domain, the Guaduas For-mation shows its maximum thickness in the cen-tral area of Checua (1098 m), and decreases to lessthan 600 m in northwards, southwards and east-wards directions (Table 1; Fig. 2a). Lower and mid-dle intervals of the Guaduas Formation range inage from uppermost Campanian to Maastrichtian,and only the uppermost strata yielded a lowerPalaeocene age (Fig. 2b). The Cacho Formationyielded four detrital zircon ages in the range of61–64 Ma in the Checua area, yielding an earlyPalaeocene age of deposition (Fig. 2b). Thickness

variations are similar to the Guaduas Formation,with a maximum thickness of 245 m in the Checuaarea. The stratigraphical thickness of the BogotaFormation is very variable, with a maximum of1415 m in Usme and a minimum of 169 m inTunja; regionally, its thickness decreases north-wards. Detrital volcanic zircon ages and palynologyindicate a middle Palaeocene–lower Eocene agefor this unit. The middle Eocene Regadera Forma-tion disconformably overlies strata of the BogotaFormation (Bayona et al. 2010).

In the eastern axial domain, the thickness ofthe Guaduas Formation varies from 290 to 609 m(Table 1; Fig. 2a) and the upper interval of theGuaduas Formation is Maastrichtian in age (Fig.2b). In Umbita only, an eastwards decrease inthickness is documented from 525 to 350 m. Thestratigraphical thickness of the Lower Socha For-mation decreases northwards from 257 m inUmbita to a range between 90 and 138 m in Pazde Rıo (Pardo 2004; Saylor et al. 2011). Similarly,the thickness of the Upper Socha decreases from450–460 m in Umbita and San Antonio to 255–280 m in Pesca and Paz de Rıo. Saylor et al. (2011)documented an abrupt change from 227 to 344 min the Upper Socha Formation over a distance ofless than 3 km. The Lower Socha Formation unitis poorly dated as lower–middle Palaeocene in thePaz de Rıo area (Pardo 2004) (Fig. 1b), whereasthe Upper Socha Formation has palynological andmineral-age constraints indicating a late Palaeo-cene–early Eocene age (Table 1; Fig. 2b). The over-lying Picacho Formation has been assigned anearly–middle Eocene age (Pardo 2004).

The Guaduas Formation in the eastern domainis less than 50 m thick and yielded only a Maastri-chtian age (Parra et al. 2009a); coeval strata in theCusiana area are not reported (Fig. 2a). This unitoverlies upper Campanian–lower Maastrichtianstrata of the Guadalupe Formation (Guerrero &Sarmiento 1996). The sandy Barco Formation isearly–middle Palaeocene in age, whereas the Cuer-vos Formation has palynological and mineral-agedata that indicate a late Palaeocene–early Eoceneage (Bayona et al. 2008; Parra et al. 2009a)(Table 1; Fig. 2). The thickness of both the Barcoand Cuervos units decreases eastwards, as wellas the thickness of the lower–middle EoceneMirador Formation (Parra et al. 2009a).

Chronostratigraphical correlation

Because stratigraphical units are diachronous(Fig. 2b), we established the following boundariesat 65, 59, 55 and 49 Ma for chronostratigraphicalcorrelation of strata. The following criteria wereused to determine the position of those times for



Table 1. Measured stratigraphical thickness, palynological results and maximum age of deposition for Maastrichtian–Eocene strata (data from Bayona et al. 2012;error for the youngest age of a single crystal is not shown). Italic font correspond to the oldest unit studied in each area. These areas correspond to the location ofmajor syncline structures, as shown in Figure 1

Lithologicalunit

Thickness (m) Palynology Mineral age(number of zircons

for mean calculation)Westernflank

Easternflank

No. ofsamples

Results Reworked pollen

Western domainGuaduero (this study) Upper Hoyon 230 220 3 Base: Middle–Late Eocene Campanian–

MaastrichtianMiddle Hoyon 0 0Lower Hoyon 0 0Seca 490 .703 22 Base: Campanian–

Maastrichtian boundaryTop: 61.9 +2 Ma (5);

60 Ma (youngestage)

Umir/Cimarrona 2 Top: Late CampanianSan Juan Rıo Seco; western

flank: Gomez et al. (2003);eastern flank (this study)

Upper Hoyon 460 250 19 Top: Middle Eocene–EarlyOligocene

Late Cretaceous

Middle Hoyon 310 1185 13 No recoveryLower Hoyon 667 595 11 No recoverySeca 400 .470 3 No recoveryUmir–Cimarrona 350 4 Top: Campanian–

Maastrichtian boundary

Western axial domainPaipa (this study) Bogota 485 1 No recovery

Cacho 200 5 No recoveryGuaduas 650 12 Top: Early Palaeocene;

metre 500 and 600is 65 Ma

Tunja (this study) Bogota 169 11 Top: Palaeocene–Eocene Campanian–Maastrichtian

Cacho 155 3 No recoveryGuaduas 600 14 Top: Late Maastrichtian;

Middle: Campanian–Maastrichtian boundary

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Checua (western flank(Sarmiento (1992 and thisstudy); eastern flank (thisstudy)

Bogota .629 .484 33 Top: Early Eocene–Oligocene;Base: middle–latePalaeocene

Top: 58 Ma (youngestage)

Cacho 219 245 6 No recovery Bottom:64.3 + 2.1 Ma (3)

Guaduas 908–1098 60 Top: Early Palaeocene;65 Ma at metre 715;Middle: Maastrichtian;Base: Campanian–Maastrichtian

Macheta (this study; mineral agefrom Horton et al. 2010a, b)

Bogota 646 7 Top: Early Eocene EarlyPalaeocene–LateMaastrichtian

Top: 58.4 + 1.2 (3)

Cacho 84Guaduas 669

Guatavita (this study) Bogota 883 1180 5 Top: Early EoceneCacho 170 4 No recoveryGuaduas .450 3 No recovery

Usme (Bayona et al. 2010) Bogota 1415 Top: Early Eocene;Middle: Base: middle–late Palaeocene;Base: Palaeocene?

Campanian–Maastrichtian

In metre 530:56.2 + 1.6 (6)

Cacho 100 Palaeocene?Guaduas .600 Top: Palaeocene?

Eastern axial domainPaz de Rıo (thickness and age

from Pardo 2004; mineral agefrom Saylor et al. 2011, 2012)

Upper Socha 280 Top: Early Eocene; Middleand base: LatePalaeocene

Top: 55.4 + 0.5 (6);Middle: 58.2+2 (8)

Lower Socha 138 Early–middle PalaeoceneGuaduas 400 Maastrichtian

Pesca (this study) Upper Socha 255 8 Base: middle–latePalaeocene

Lower Socha 125Guaduas 290 7 Middle and Base: Late

Campanian–LateMaastrichtian

(Continued)

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Table 1. Continued

Lithologicalunit

Thickness (m) Palynology Mineral age(number of zircons

for mean calculation)Westernflank

Easternflank

No. ofsamples

Results Reworked pollen

Rondon (this study) Upper Socha 355 25 Top: Early Eocene; 55 Mais between metres 100and 200

Lower Socha 200 4 Middle Palaeocene? LateMaastrichtian

Guaduas 295 61 Top: Late Maastrichtian;Base: EarlyMaastrichtian

San Antonio (this study) Upper Socha 324–460 6 No recoveryLower Socha 225–262 No recoveryGuaduas 609 3 Latest Campanian–

MaastrichtianUmbita (this study) Upper Socha 450 400 5 Base: middle–late

PalaeoceneMaastrichtian Top: 53 Ma (youngest

age)Lower Socha 230 257 2 Early–middle PalaeoceneGuaduas 525 350 19 Base: Maastrichtian

Eastern domainCusiana (Bayona et al. (2008)) Mirador 141

Cuervos 158Barco 56Guaduas 0

Medina (western flank fromParra et al. 2009a); easternflank (from this study and adeep well)

Mirador 237 135 Early–middle EoceneCuervos 477 348–440 Middle–late Palaeocene 55.5 + 0.6 (9);

57.6 + 0.7 (6)Barco 231 125–187 Early–middle PalaeoceneGuaduas 50 53 Maastrichtian

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each stratigraphical section (Fig. 3) and the calcu-lation of sedimentation rates presented in Table 2.

† Palynological data and maximum ages of depo-sition (Table 1) were used as the most reliableinformation for age constraints.

† When the stratigraphical unit does not have anage, we used the age determined for the same unitin an area within the same structural domain.

† Sandy units (i.e. amalgamated channels in theCacho Formation) are considered to have lowerrates of accumulation than muddy units (i.e.the floodplains in the Bogota Formation).

Campanian–lower Maastrichtian conglomerate-bearing strata are at the western and easterndomains, while fine-grained strata of the GuaduasFormation accumulated over sandstones of theGuadalupe Formation in the axial zone (Fig. 2b).Onset of fine-grained deposition of the Guaduasand Seca formations is older in the western domainthan in the axial domain; in this later domain,onset of deposition is nearly coeval (Fig. 2b). Thelacuna (non-deposition + erosion) overlying Secaand Guaduas strata increases eastwards betweendomains, and northwards in the four domains. Onsetof fine-grained deposition of the Bogota–Cuervos

Fig. 3. Identification of chronostratigraphical markers in the stratigraphical column of the Umbita area, and the locationof boundaries of chronostratigraphical intervals. Note the change in rate of accumulation between the sandy andfine-grained intervals. See Table 2 for the identification of chronostratigraphical intervals in other areas.



Table 2. Estimated stratigraphical thickness for chronostratigraphical intervals in each study area (see Fig. 3 for an explanation of the methods) and the calculationof sedimentation rates (m Ma21) shown in Figure 8. Italics and roman fonts correspond to the interval of 65–70 Ma in each area. Numbers in parentheses in thelithological units column make reference to stratigraphic thickness in metres

Domain Chronostratigraphicalinterval

(Ma)

Western flank Lithologicalunits

Eastern flank Lithologicalunits

Thickness(m)

Rate(m Ma21)

Thickness(m)

Rate(m Ma21)

Western domainGuaduero 49–55 20 3.3 Seca 74 12.3 Seca

55–59 38 9.5 Seca 63 15.8 Seca59–65 352 58.7 Seca 456 76.0 Seca65–70 17 3.4 Seca no data –

San Juan Rio Seco 49–55 317 52.8 Middle Hoyon 833 138.8 Middle Hoyon55–59 442 110.5 Lower Hoyon 610 152.5 Lower Hoyon (260);

Middle Hoyon (350)59–65 472 78.7 Seca (272); Lower

Hoyon (200)733 122.2 Seca (397); Lower

Hoyon (336)

65–70 478 95.6 Cimarrona (347);Seca (131)

no data –

Western axial domainPaipa 49–55 160 26.7 Bogota

55–59 325 81.3 Bogota59–65 322 53.7 Guaduas (122);

Cacho (200)

65–70 520 104.0 GuaduasTunja 49–55 70 11.7 Bogota

55–59 130 32.5 Cacho (30); Bogota (100)59–65 224 37.3 Guaduas (100);

Cacho (124)

65–70 500 100.0 GuaduasChecua 49–55 278 46.3 Bogota

55–59 575 143.8 Bogota59–65 498 83.0 Guaduas (253);

Cacho (245)

65–70 655 131.0 Guaduas

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Macheta 49–55 278 46.3 Bogota55–59 368 92.0 Bogota59–65 334 55.7 Guaduas (250);

Cacho (84)

65–70 419 83.8 GuaduasGuatavita 49–55 708 118.0 Bogota

55–59 472 118.0 Bogota59–65 340 56.7 Guaduas (170);

Cacho (170)

65–70 280 56.0 GuaduasUsme 49–55 693 115.5 Bogota

55–59 526 131.5 Bogota59–65 325 54.2 Guaduas (225);

Cacho (100)

65–70 375 75.0 Guaduas

Eastern axial domainPaz de Rio 49–55 90 15.0 Upper Socha (44);

Picacho (46)55–59 188 47.0 Upper Socha59–65 185 30.8 Lower Socha (138);

Upper Socha (47)

65–70 400 80.0 GuaduasPesca 49–55 130 21.7 Upper Socha (73);

Picacho (57)55–59 178 44.5 Upper Socha59–65 125 20.8 Lower Socha

65–70 232 46.4 GuaduasRondon 49–55 163 27.2 Upper Socha

(130);Picacho (33)

55–59 265 66.3 Lower Socha (40);Upper Socha(225)

59–65 160 26.7 Lower Socha

65–70 236 47.2 GuaduasSan Antonio 49–55 291 48.5 Upper Socha (225);

Picacho (66)

(Continued)

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Table 2. Continued

Domain Chronostratigraphicalinterval

(Ma)

Western flank Lithologicalunits

Eastern flank Lithologicalunits

Thickness(m)

Rate(m Ma21)

Thickness(m)

Rate(m Ma21)

55–59 280 70.0 Lower Socha (40);Upper Socha (240)

59–65 219 36.5 Lower Socha

65–70 487 97.4 GuaduasUmbita 49–55 250 41.7 Upper Socha (196);

Picacho (54)55–59 245 61.3 Lower Socha (41);

Upper Socha (204)59–65 214 35.7 L Socha

65–70 280 56.0 GuaduasEastern domainCusiana 49–55 54 9.0 Mirador

55–59 128 32.0 Cuervos59–65 86 14.3 Barco (56);

Cuervos (30)

65–70 0 0.0Medina 49–55 96 16.0 Mirador 0 0.0

55–59 477 119.3 Cuervos 261 65.3 Cuervos59–65 231 38.5 Barco 274 45.7 Barco (187);

Cuervos (87)

65–70 50 10.0 Guaduas 53 10.6 Guaduas

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strata is coeval in the eastern and axial domains,whereas, in the western domain, onset of fine-grained deposition of the middle Hoyon Formationis younger than in the axial domain.

The Palaeocene–Eocene boundary is well con-strained by palynology and detrital zircon ages inthe four domains (Fig. 2b), showing a correlation,from west to east, of the middle Hoyon–Bogota–Upper Socha–Cuervos formations (Fig. 2b). LowerEocene strata in the eastern domain and easternaxial domain include sandstones of the Mirador–Picacho formations, whereas, in the western axialdomain, it corresponds to the fine-grained strataof the upper Bogota Formation. In the westerndomain, the middle Hoyon Formation in the southand strata mapped as upper Seca Formation in thenorth have an inferred mid-Palaeocene–lowerEocene age (Fig. 2b).

Provenance

The description of provenance signatures (sand-stone petrography, heavy minerals and U–Pbdetrital zircon geochronology) follows the chronos-tratigraphical intervals defined earlier to understandthe dispersal of terrigenous detritus by intervalof time (see Figs 4–6). Glauconite was identifiedeither as authigenic (e.g. clean faces filling pores:Fig. 4e) or as reworked fragments (clay coatingand rounded fragments: Fig. 4f, g).

Upper Campanian sandstones (.70 Ma)

Quartzarenites show authigenic glauconite, as ident-ified at the top of the Guadalupe Formation in theUmbita (Fig. 4e) and Checua areas (e.g. Sarmiento1992). Heavy mineral assemblage of the GuadalupeFormation in the eastern domain is dominantlyultrastable (Fig. 5a) but the presence of titanitemakes up 11.7% of the sample (Fig. 5b).

Detrital zircons for Campanian strata from theaxial and eastern domains reveal a dominant Pro-terozoic population, with several age peaks betweenapproximately 0.96 and 1.8 Ga (see also the Cre-taceous rock data in Horton et al. 2010b andSaylor et al. 2011), whereas one sample from thewestern domain additionally reveals a youngerpopulation (,0.3 Ga) (Fig. 6). In these three areas,Palaeozoic age populations are a minor population.

Maastrichtian sandstones (65–70 Ma)

The provenance signature of Maastrichtian sand-stones in the axial and eastern domains fall towardsthe field of craton interior, whereas sandstonesof the western domain fall in the field of transi-tional recycled orogen (Fig. 4a). However, lithic

fragments are dominantly sedimentary in all of thefour domains. Reworked glauconite was identifiedin five areas of the axial domain (Checua, Paipa,Umbita, Rondon and Pesca) (Fig. 4f, g). Heavymineral assemblage is dominantly ultrastable (Fig.5a), with stable minerals in the western domain(apatite and garnet, Guaduero area: Fig. 5c) andenrichment of micas in northern areas of the axialzone (Rondon and Paipa).

The northern segments of the western axialand eastern axial domains are characterized by theappearance of Phanerozoic detrital zircon popu-lations, including Cretaceous (c. 70–80 Ma) andSilurian–Cambrian (c. 424–534 Ma), as well asProterozoic zircons similar to the population ofCretaceous strata (Figs 6 & 7). Cretaceous ages(65–100 Ma) are the dominant age population inthe western domain. In the Umbita area, the agepopulation is older than 1.2 Ga, and a single Guad-uas Formation sample that shows ages older than0.9 Ga (Saylor et al. 2011).

Lower–middle Palaeocene sandstones

(59–65 Ma)

The provenance signature defined by the sandstonepetrography of the eastern and western domainsvaries in the content of the amount of quartz andlithic fragments (craton interior and transitionalrecycled orogen, respectively), whereas provenancesignature of sandstones of the two axial domainsoverlap in the field of recycled quartzose orogen(Fig. 4b). Total lithic fragment content increaseswestwards (Fig. 4b, h–j). Metamorphic and pluto-nic rock fragments increase with respect to Maas-trichtian sandstones; these fragments, volcaniclithic fragments and feldspars are more abundant inthe western domain and the eastern axial domain(Fig. 4b). Reworked glauconite fragments wereidentified in the Checua and Usme areas of thewestern axial domain. Unstable and stable heavymineral assemblages are more abundant in thewestern domain and in the Usme area, whereas, inthe other areas of the axial zone and the easterndomain, ultrastable minerals dominate (Fig. 5a).Heavy minerals concentrate in laminae in theUmbita area, and in Pesca and Paipa muscovite isvery common. Garnets fragments were identifiedin the Umbita area (eastern axial domain: Fig. 5d),while pyroxene fragments were identified in SanJuan de Rıo Seco (western domain: Fig. 5e) andUmbita (eastern axial domain: Fig. 5f ).

Cretaceous detrital zircons, together with morelimited Triassic (248–230 Ma), Early Palaeozoicand Proterozoic detrital zircons, dominate in thewestern domain and the Usme area (Figs 6 & 7),as well as in the Paz de Rıo area (eastern axial



Fig. 4. (a)–(d) Ternary provenance and lithical population diagrams for each chronostratigraphical interval, showingthe calculated mean value. The shaded polygons represent 1s uncertainty envelopes. For lower Eocene sandstones(Fig. 4d), samples of the quartzose Picacho and Mirador formations are plotted separately because those units are abovean unconformity. Note the upsection increase in lithical content in all domains, and the enrichment of metamorphiclithical fragments in late Palaeocene time for the western and eastern domains. (e) Authigenic glauconite. (f ) & (g)Reworked glauconite and ferruginous matrix reported in Maastrichtian strata of the axial domain. Areas with reworkedglauconite are indicated for each interval of time. (h)–( j) Eastwards increase in sandstone maturity of lower–middlePalaeocene sandstones, indicating at least three provenance fields (see b) from transitional recycled orogen to the west,to quartzose recycled orogen in the axial zone, and to craton interior to the east.



domain: Fig. 6) (Saylor et al. 2011). Although Pha-nerozoic zircon populations are similar along thewestern axial zone, there is a northwards increaseof Proterozoic (0.8–2 Ga) age populations (Fig.7). In contrast, one sample from the eastern domainshows older zircon age peaks of 600, 900 and1500 Ma (Fig. 6).

Upper Palaeocene sandstones (55–59 Ma)

The provenance signature of the four studieddomains plot near the boundary of quartzose andtransitional recycled orogens, being more lithic

and feldspar-rich in the western domain (Fig. 4c).The content of total lithic fragments increasesslightly with respect to lower–middle Palaeocenesandstones. Metamorphic–plutonic rock fragments,as well as feldspar fragments, are more abundantin western, eastern and western axial domains,whereas sedimentary rock fragments dominate inthe eastern axial domain (Fig. 4c). Reworked glau-conite was identified in the Usme (western axialdomain) and Umbita (eastern axial domain) areas.Unstable heavy minerals were found in the Usme(epidote), Rondon (chlorite) and Medina (hornble-nde: Fig. 5g) areas, and in the former two areasthe unstable population is an important constituent

Fig. 5. (a) Lateral distribution of heavy minerals in Maastrichtian–lower Eocene sandstones. (b)–(h)Photomicrographs of titanite, garnet, pyroxene and hornblende as examples of unstable and stable minerals reported atdifferent domains and at different age intervals; the presence and good preservation of these minerals indicateprovenance from metamorphic and igneous rocks, and accumulation in a first-cycle sandstone. (i) Muscovitefragments are very common in northern areas of the axial zone; this mineral is an indicator of accumulation inlow-energy settings.



(Fig. 5a). In the other areas of the western, axial andeastern domains, the ultrastable assemblage domi-nates (Fig. 5a). Micas continue as important con-stituents in northern areas (Paipa–Pesca) of theaxial zone. Garnet was identified in the Guaduero(western domain) and Medina (eastern domain)areas.

The detrital zircon age populations from the west-ern domain are dominantly Mesozoic (90–250 Ma:Fig. 6), whereas samples from the western axialand eastern axial domains show an increase in Pha-nerozoic detrital zircon ages (65–360 Ma) anddecrease in Proterozoic (.1000 Ma) zircon ageswhen compared with the lower–middle Palaeocene

Fig. 6. Detrital zircon age population distributed by age interval and structural domain. Note the dominance of theyoung age population (65–360 Ma) in the western domain since the Campanian, while this population is absent in theeastern domain. Only Palaeocene ages of volcanic zircons are reported in the eastern domain. In the western axial zone,the dominance of 65–360 Ma increases in lower–middle Palaeocene sandstones, whereas, in the eastern axial zone, thispopulation increases in upper Palaeocene sandstones. Palaeozoic and Proterozoic age populations dominate in theeastern axial and eastern domains in all of the intervals, and in the western axial domain in upper Cretaceous rocks,and lower–middle Palaeocene rocks of the northern areas (see Fig. 7). Archean ages (.2500 Ma) are absent in all of thestudy areas.



samples (Fig. 6). In contrast, the eastern domainshows a bimodal distribution with a minor popu-lation of 50–65 Ma zircons and a prominent Proter-ozoic population (.1.5 Ga) (Fig. 6).

Lower Eocene sandstones (49–55 Ma)

The provenance signature of the Mirador (easterndomain) and Picacho (eastern axial domain)

sandstones plot in the upper boundary of quartzoserecycled orogen, whereas the other sandstones ofthe western axial and eastern axial domains(Bogota–Upper Socha formations) plot near theboundary of quartzose and transitional recycledorogens (Fig. 4d). The former two units are abovethe unconformity, while the latter two units arebelow the unconformity. The content of metamor-phic–plutonic and sedimentary rock fragments is

Fig. 7. Detrital zircon age populations of lower–middle Palaeocene sandstones from the western axial zone, showingthe northwards increase in 900–1800 Ma ages. The increment of this age population is related to the reworking ofCretaceous rocks, as the exposure of these units increases northwards (see Fig. 1).



proportional in these sandstones, with an increase inthe content of the latter fragments in the Picachosandstones (Fig. 4d). The highest content of feld-spar fragments is in the western axial domain.Reworked glauconite was identified in the Usme(western axial domain) and Umbita (eastern axialdomain) areas. Stable and unstable heavy mineralassemblages are common both in the Usme (garnetin Fig. 5h: epidote and hornblende) and Umbitaareas (garnet, titanite, epidote and hornblende),with a dominance of micas in the Rondon area(Fig. 5a, i). The ultrastable assemblage becomesdominant in quartzose sandstones of the Picachoand Mirador formations, above the unconformity(Fig. 5a); however, in the Medina area (easterndomain), apatite, hornblende, pyroxene and titaniteare present at levels of between 2 and 5%.

Detrital zircon age populations from the west-ern axial domain mostly include those of less than290 Ma, with significant 58–70 and 220–290 Maage peaks. Whereas, in the eastern axial domain,similar Phanerozoic detrital zircon age popula-tions alternate with more abundant Proterozoic(.1000 Ma) zircon ages, with peak ages as old asabout 1.8 Ga (Fig. 6). Samples from the UpperSocha Formation, reported in Saylor et al. (2011),include zircons in the range of 53–57 Ma. In con-trast, detrital zircon populations reported for onesample of the Mirador Formation in the easterndomain (Horton et al. 2010a, b) is characterizedby zircons older than 1.4 Ga.

Spatial and temporal changes in

sedimentation rates

Division of stratigraphical units (thickness andlithologies) into chronostratigraphical intervals(age) enable sedimentation rates to be calculated(see Table 2 for a summary). In this section, wedocument the spatial trend of sedimentation ratesfor each interval of time in order to locate depo-centres, and infer its three-dimensional geometryin map and profile view (Figs 8 & 9). We use asedimentation rate of 80 m Ma21 as an arbitraryreference to enclose areas with relatively high sedi-mentation rates as depocentres.

Maastrichtian depocentres (65–70 Ma)

Two depocentres with localized areas of maxi-mum sedimentation rates were formed at thistime. In the western domain, a localized depo-centre formed around the San de Juan de Rıo Secoarea (95 m Ma21). Sedimentation rates decreaseabruptly further north in the Guaduero area(3.4 m Ma21), and we infer that sedimentationrates of the depocentre increased eastwards (i.e.

become thicker eastwards: Fig. 9b). Maastrichtianstrata thin or are eroded by an unconformity fur-ther to the north (Rıo Minero area: Restrepo-Paceet al. 2004) and are not recorded by an unconfor-mity further south (Fusa Syncline: Bayona et al.2003). In contrast, the depocentre in the axialdomain has an elongated geometry with a NNWstrike (Fig. 8a). A maximum sedimentation rate of131 m Ma21 is documented in the Checua area;other values decrease in all directions as far as theeastern domain, where Maastrichtian strata are notpreserved (Fig. 8a). The profile geometry of theaxial depocentre is like asymmetrical sag that isthicker to the west.

Lower–middle Palaeocene depocentres

(59–65 Ma)

Maximum sedimentation rates of 122 m Ma21 inthe western domain are located in the East SanJuan de Rıo Seco region and decrease northwardsto the Guaduero area (58–76 m Ma21), suggestinga north-striking elongated depocentre geometry(Fig. 8b) that is thicker eastwards (Fig. 9c). In con-trast, the former Maastrichtian depocentre in theaxial domain broke into two depocentres. In thewestern axial domain, the Checua area continuesas the area with a maximum sedimentation rate of83 m Ma21; sedimentation rates decreases slightlysouthwards to 55–57 m Ma21 and more abruptlytowards the eastern axial domain, where sedimen-tation rates are between 30 and 36 m Ma21. In theeastern domain, a new depocentre began to form assedimentation rates increased to 38–45 m Ma21.The profile geometry of the axial depocentre con-tinues as an asymmetrical sag that is thicker west-wards, whereas the geometry of the easterndomain depocentre has a wedge-like geometry thatis also thicker westwards (Fig. 9c).

Upper Palaeocene depocentres (55–59 Ma)

During late Palaeocene time, the depocentre in thewestern domain had a similar pattern to that descri-bed before, with maximum sedimentation rates inthe East San Juan de Rio Seco area (152 m Ma21)and a profile geometry that is thicker eastwards(Fig. 9c). The other two depocentres described forthe axial and eastern domains, respectively, have asignificant increase in sedimentation rates in com-parison with rates reported in lower–middle Palae-ocene time (Fig. 8c), but the profile geometryremains similar. In the axial domain, sedimentationrates increase to 92–143 m Ma21 in the Checua–Guatavita areas, whereas, further south, they reached131 m Ma21 in the Usme area. In other areas furthernorth of the axial domain, sedimentation rates



Fig. 8. (a)– (d) Palaeogeographical maps for the four chronostratigraphical intervals discussed in the text, showingthe location of source areas, active structures, geometry of depocentres, sedimentation rates, depositionalenvironments and dispersion of terrigenous detritus. (e) Palinspastic map for the Palaeogene time ofSarmiento-Rojas (2001), used as base map for the location of studied areas shown in (a)–(d). These diagrams show thenorthwards withdrawal of marginal deposits and how depositional systems of the western domain have beendisconnected since Maastrichtian time. (b) & (c) During the Palaeocene, a single depositional system broke into twodepositional systems due to eastwards migration of intraplate reactivation. The depositional system in the axial zoneshows the presence of low-energy settings (i.e. continental lakes); further north, this depositional system is closed by theuplift of the Santander Massif at this time (Ayala et al. 2012). In contrast, the depocentre formed in the eastern domainshows northwards dispersion of terrigenous material. Depocentres with sedimentation rates in excess of 80 m Ma21 inthe axial domain formed eastwards of reactivated faults. (d) In early Eocene time, magmatic activity continues along theCentral Cordillera and advances onwards to intraplate settings (Fig. 9c) (Bayona et al. 2012); localized depocentresformed in the southern western axial domain and in the western domain, and a low tectonic subsidence regimedominated in the eastern axial and eastern domains.



Fig. 9. (a)–(c) Crustal-scale cross-sections for 75, 65 and 55 Ma (see Fig. 8a, c for the location) showing therelationship between Caribbean subduction, marginal and intraplate magmatism (from Bayona et al. 2012), crustaltilting, fault reactivation, and geometry of intraplate syn-orogenic basins. Tectonic subsidence mechanisms varyeastwards from crustal tilting of the Central Cordillera to reactivation of intraplate structures. Both mechanisms arerelated to subduction of the Caribbean Plate. In early Eocene time, reactivation in intraplate structures in both margins ofthe Eastern Cordillera formed a sag-like basin bordered by exposed Cretaceous rocks; deposition in the Llanos Basincorresponds to filling of a foreland basin (Bayona et al. 2008). (d) & (e) Regional tectonic setting, showing the shift inthe style of collision of the Caribbean Plate with the NW continental margin, and its effects on continental arcmagmatism and reactivation of intraplate structures (modified from Bayona et al. 2012).



range between 47 and 81 m Ma21. In the depocentreof the eastern domain, sedimentation rates reached119 m Ma21 for the Medina area.

Lower Eocene depocentres (49–55 Ma)

Only two depocentres formed at this time. Thedepocentre in the western domain maintains itslocation in the San Juan de Rio Seco area (up to138 m Ma21), with a wedge-like geometry thicken-ing eastwards. As in previous intervals of time, sedi-mentation rates decrease abruptly northwards toless than 12 m Ma21. The other depocentre formedin the southern western axial domain, in the Usme(115 m Ma21) and Guatavita (118 m Ma21) areas.Further north and east, sedimentation rates decreaseto values ranging from 15 to 45 m Ma21 in the axialdomain and from 0 to 16 m Ma21 in the easterndomain.

Discussion

Integration of provenance data, spatial and temporalvariations in sedimentation rates, and location ofdepocentres, using the systematic analysis carriedout along the 16 studied sections (Fig. 1c), allowsus to propose a new model for the tectonostratigra-phical evolution of Maastrichtian–lower Eocenesyn-orogenic basins, and to define how these syn-orogenic basins formed by the interaction of theCaribbean and South American plates. In thispaper, we compare our model with pre-existing tec-tonostratigraphical models, suggesting either fore-land basin setting or negative flexural basin for theearly Palaeogene period.

Maastrichtian–early Eocene

tectonostratigraphical model and its

relationship to Caribbean tectonics

Provenance signatures of Maastrichtian–middlePalaeocene sandstones (Fig. 4a, b) show at leasttwo different source areas supplying sediments tothe western and the axial domains. Supply frombasement rocks of the Central Cordillera is welldocumented in both domains by the presence ofa 65–360 Ma detrital zircon population, metamor-phic lithic fragments, as well as stable and unstableheavy mineral associations. The emergence ofthe Central Cordillera in latest Cretaceous–earlyEocene has been also documented using thermo-chronological and provenance data (Gomez et al.2003; Villagomez et al. 2011; Caballero et al.2013).

The other source area should be represented bythe Cretaceous sedimentary cover separating the

western domain from the axial domain, as indicatedby the following criteria.

† The presence of reworked glauconite (Fig. 4f, g)and Cretaceous pollen in different areas of theaxial domain (Fig. 8a, b).

† An increase in muddy matrix in the fabric ofPalaeocene sandstones (Fig. 4f, g).

† A contrasting upsection change in heavy mineralassemblages. In the western domain, stable andunstable minerals are more common in lower–middle Palaeocene sandstones, and ultrastableminerals become dominant in upper Palaeocenesandstones (Fig. 5a). In the axial domain, stableand ultrastable heavy mineral associations dom-inate with a localized increase in unstable min-erals in the Usme area in upper Palaeocenesandstones.

† A northward enrichment in 900–1800 Ma detri-tal zircon age population in sandstones of theaxial zone (Fig. 7).

The criteria presented above also indicate that theuplifted area must have been located close to theplace of deposition to the axial sections, as pollenand glauconite are very unstable fragments tosurvive tropical weathering and more than 40 kmof fluvial transport, as documented in the fluvialsands of the Llanos Basin of Colombia (Amorochoet al. 2011). The reworking of compositionallymature Cretaceous sandstones concentrates ultra-stable heavy minerals. The muddy fraction may bea product of the erosion of palaeosols (Potter et al.2005), and 900–1800 Ma detrital zircon ages are avery common population in Cretaceous rocks(Horton et al. 2010a, b).

Separation of the western domain from the axialdomain by the uplift of the Cretaceous sedimentarycover is also suggested by the eastwards increasein thickness in the western domain, and high sedi-mentation rates reported in the San Juan de RioSeco area. Tectonic loads eastwards of the Gua-duero and San Juan de Rıo Seco sections favouredhigher tectonic subsidence of the western domain(Fig. 9a, b). Those tectonic loads correspond to theinversion of former normal faults of the westernmargin of the former extensional basin (Figs 8a &9b), as proposed by Sarmiento-Rojas (2001) andBayona et al. (2008), who used geodynamic models,and Cortes et al. (2006), who used the surface andsubsurface mapping relationships of Palaeogeneunits (Fig. 1).

Uplift of the western flank of the Eastern Cor-dillera was not continuous southwards becausecontinental deposits in the western domain fur-ther south of the San de Rio Seco area were con-nected with a fluvial system in the Usme area, assuggested by similar detrital zircon populations



between these two areas (Figs 6 & 8a, b). North-wards and eastwards dispersion of detritus sup-plied from the Central Cordillera in the axialdomain is further supported by palaeocurrent indi-cators reported for Palaeocene sandstones in theaxial domain (see Bayona et al. 2008 for a reviewand Saylor et al. 2011 for palaeocurrent data in thePaz de Rıo area).

Reactivation of the western margin of theformer extensional basin may be also explainedby crustal tilting of the Central Cordillera–SantaMarta Massif (Fig. 9a, b), an uplift mechanismof the continental margin related to the subduc-tion of the Caribbean Plate (Bayona et al. 2011,2012; Cardona et al. 2011; Ayala et al. 2012).The presence of: (1) metamorphic rock fragments;(2) garnet, apatite and epidote fragments; and (3)the dominance of Cretaceous and Permo-Triassiczircon age populations over Jurassic age popula-tions support the erosion of Cretaceous intru-sive rocks and Permo-Triassic metamorphic rockspresently exposed adjacent to the Romeral palaeo-suture.

Reactivation of the western fault system ofthe former extensional basin acted as the easternboundary of the tilted rigid crustal block (Fig. 9).Eastwards thickening of the depocentre in thewestern domain (Figs 8 & 9) is also explained byeastwards tilting of the Central Cordillera. Simi-larly, eastwards tilting of the inverted hanging-wallblock (Figs 9b, c) explains the elongated depocen-tre documented in Maastrichtian–late Palaeocenetime (Fig. 8a–c) in the western axial domain. Sedi-mentation rates in the western axial depocentreare equivalent to, or higher than, the rate determinedin the eastern San Juan de Rio Seco area in thewestern domain. The simultaneous eastwards-tilting activity of uplifted blocks may explain thesimilar sedimentation rates in two areas that are130 km apart (Fig. 8). However, local structure-controlled patterns of sedimentation in the westerndomain (e.g. normal faulting in Fig. 8c) controlledlocalized deposition of fine-grained strata of themiddle Hoyon Formation.

The distance between the Central Cordillera andthe eastern axial and eastern domains is greaterthan 250 km (Fig. 8). Therefore, chemically unsta-ble fragments are not expected to survive hundredsof kilometres of fluvial transport in tropical set-tings (Johnsson et al. 1991; Amorocho et al. 2011).Reworked Cretaceous pollen, unstable heavy min-eral fragments (Fig. 5: garnet, titanite, hornblende,epidote and chlorite) and a high content of lithicfragments are common features in upper Palaeo-cene–lower Eocene sandstones of the eastern axialand eastern domains (Fig. 4c, d). Therefore, unsta-ble terrigenous detritus in the eastern axial andeastern domains were not entirely derived from

the Central Cordillera. In addition, a characteristicdetrital zircon population derived from the CentralCordillera (65–360 Ma) is not recorded in theeastern domain (Fig. 6).

In the late Palaeocene, mud-dominated depo-sition and an increase in the content of metamorp-hic lithic fragments in sandstones of the easterndomain began earlier than in the axial domain(Figs 2b & 4c). Unstable titanite and hornblendefragments in Campanian and upper Palaeocenestrata (Fig. 5a) also indicate the presence of nearbylocated basement-cored uplifts. The detrital zirconage population of the upper Palaeocene strata inthe eastern domain (Fig. 6) is similar to the zirconpopulation older than 1.2 Ga reported from a bur-ied basement-cored high in the proximal LlanosBasin (Ibanez-Mejia et al. 2009). These buriedbasement-cored highs, presently under the Cenozoicforeland strata, include Cretaceous and/or Palaeo-zoic sedimentary units overlying basement rocks(Bayona et al. 2007). Therefore, erosion of thesedimentary cover and basement-cored uplifts sup-plied quartzose detritus, metamorphic lithic frag-ments, unstable heavy minerals and detritalzircons older than 1.2 Ga.

Reactivation of the eastern normal fault sys-tem of the former extensional basin and the devel-opment of structures in the southern Llanos Basinoccurred during the Palaeocene (see Mora et al.2013). This deformation is evident in late Palaeo-cene time by the immature composition of sand-stones (Fig. 4c), and sedimentation rates in theeastern domain are higher than those in the easternaxial domain (Figs 8c & 9c). The location of thedepositional axis in the eastern domain during theMaastrichtian (Figs 8a & 9b) and the reactiva-tion of the eastern normal fault system precludethe eastwards transport of detritus supplied fromthe Central Cordillera to the eastern domain (Fig.6). Tectonic loading between the western and east-ern domains contributed to the flexure-driven sub-sidence and foreland pattern of the depositionbetween the eastern axial domain and the proximalLlanos Basin in Maastrichtian–late Palaeocenetime (Fig. 9b, c) (Bayona et al. 2008).

Fault reactivation along the eastern marginis coeval with the record of volcanic zircons inupper Palaeocene sandstones (Bayona et al. 2012).These two processes may be explained by shalloweastwards-dipping subduction of the CaribbeanPlate (Figs 9c, d) that produced intraplate mag-matism and fault reactivation located as far eastas the proximal Llanos Basin. Such intraplatemagmatism affecting the Eastern Cordillera andadjacent basins in the east has been reportedfor Cretaceous (Vasquez et al. 2010) and Mio-Pliocene periods (Taboada et al. 2000; Vasquezet al. 2009).



Comparison to pre-existing tectonic models

A major difference between our study and pre-existing tectonostratigraphical model studies of theNorthern Andes is the systematic analysis along13 studied sections distributed across the four struc-tural domains, and the integration of data from otherstudies along three sections. Some previous studiesfocused on data collected from only a single area(e.g. Paz de Rıo area: Saylor et al. 2011, 2012;Medina area: Parra et al. 2009a) or from a singlestructural domain (Middle Magdalena Valley:Gomez et al. 2003, 2005a, b). Other ones includeda regional study with no details of specific sections(Cooper et al. 1995; Pindell et al. 2005) or usedresults from only one technique, such as thermo-chronology (Mora et al. 2010).

The model of a continuous foreland basin model,which lasted either from Maastrichtian to middleMiocene (Cooper et al. 1995) or from Maastrichtianto late Eocene (Gomez et al. 2005a; Horton et al.2010b), predicts the following phenomena, whichare not supported by our data (see discussionabove) or by recently published data, including thefollowing.

† Burial of Upper Cretaceous rocks was homo-geneous across the western margin of theformer extensional basin. The continuous fore-land basin model implies that the thickestPalaeogene deposition occurred in the areawhere we propose the onset of inversion ofMesozoic structures (western margin of the for-mer extensional basin). Moretti et al. (2010)used vitrinite reflectance as a palaeothermome-ter to compare burial processes in the footwalland hanging wall of the Bituima Fault. Theirresults show an eastwards decrease in maturityof the Umir Formation, and they interpreted anearly Cenozoic reactivation of the Bituima–ElTrigo Fault zone, similar to the suggestion byCortes et al. (2006). Published thermochrono-metry studies indicate that exhumation of thewestern flank of the Eastern Cordillera isyounger southwards. Parra et al. (2012) andCaballero et al. (2013) document the onset ofexhumation in the Palaeocene of the westernflank of the Eastern Cordillera and northernMagdalena Basin. Exhumation ages in thehanging wall of the Bituima Fault, further tothe south, ranges from 41 to 21 Ma (Gomezet al. 2003; Parra et al. 2009b; Caballero et al.2013) and from 50 to 35 Ma for the westernlimb of the Guaduas Syncline (Parra et al.2009b). Although thermochronometry does notreveal early Palaeogene onset of cooling, prove-nance signatures and sedimentation rates record

those early Palaeogene phases of deformation,as shown in this study. In order to test the hypo-thesis proposed in our study, it is necessary tocarry out detrital thermochronology analysis inEocene and younger rocks in the Guaduas Syn-cline and Checua area.

† An east-verging thrust belt (Central Cordillera)bounding a syn-orogenic basin formed by flex-ural subsidence. In the foreland basin model,it is expected that the frontal segment of theCentral Cordillera supplied sediments to thebasin, with an expected dominance of Jurassiczircon age populations, feldspar and volcaniclithic fragments (Fig. 9a). However, heavymineral analysis and detrital zircon age popu-lations (see data above) indicate that Cretaceousintrusive rocks and Permo-Triassic metamorphicrocks, presently exposed adjacent to the Romeralpalaeo-suture, were the dominant source rocks(Fig. 9a). Therefore, the western margin of theCentral Cordillera was exposed, whereas theeastern flank of the Central Cordillera wasnot exposed at this time. Our data support themodel of an eastwards-tilting Central Cordillerawith an apparent absence of bounding thrusts(Pindell et al. 2005) as result of Caribbean sub-duction. The model proposed by Pindell et al.(2005) describes the syn-orogenic basin as a‘sag-like’ structure with a negative flexuremechanism. The irregular eastwards variationof thickness and sedimentation rates in thePalaeocene period (Fig. 8: Tables 1 & 2) sup-ports neither a positive flexural foreland basinnor a single elongated depocentre expected inthe negative flexural model.

† A single magmatic arc. Saylor et al. (2012)supported the foreland basin model for thePalaeocene time because they interpreted thatPalaeocene volcanic zircons reported in thePaz de Rio area were supplied from the Cen-tral Cordillera and transported by an east- tonortheastward-directed drainage. In this paper,we document that early Palaeogene volcanismaffected both the Central Cordillera and theintraplate settings (see Bayona et al. 2012). Wepropose that volcanic activity may be generatedas far as in the eastern domain, as indicated bythe mixture of Palaeocene volcanic zircons withzircon populations originated in the craton (Fig.6), and the presence of volcanic rocks in theUsme area. Our palaeogeographical model con-siders that Palaeocene sediments in the axialzone were supplied by the Central Cordilleraand sedimentary cover of inverted structures,following a northeastwards-directed drainagepattern, as documented by palaeocurrent mea-surements (see the data in Bayona et al. 2008;Saylor et al. 2011) (Fig. 8).



Similar provenance patterns and regular trendof sedimentation rates in the axial domain allowsthe interpretation that the Boyaca–Soapaga FaultSystem was not active during Palaeocene–earlyEocene time (Fig. 9c). Thermochronological datafrom samples in the hanging wall of those struc-tures (Mora et al. 2010; Saylor et al. 2012) docu-ment exhumation in late Eocene–Oligocene time,supporting the interpretation of a continuous basinin the axial zone. Why this fault system did notreactivate, in contrast to the faults bordering theextensional basin, needs further investigation.

Conclusions

The spatial and temporal variation of sedimenta-tion rates and provenance signatures was used todecipher early phases of deformation in a multi-phase orogen, such as the Eastern Cordillera ofColombia. Maastrichtian–lower Eocene strata inthe western, axial and eastern domains of the cen-tral segment of the Eastern Cordillera support amodel of eastwards migration of intraplate faultreactivation (positive inversion structures). Dia-chronous intraplate fault reactivation, developmentof syn-orogenic basins and intraplate magmatismare associated with the convergence and subductionof the buoyant Caribbean Plate beneath the SouthAmerican margin (Fig. 9). The reactivation offormer Cretaceous faults limited the eastwards tiltof the Central Cordillera, and compressive stressestransmitted to intraplate settings reactivated faultsas far in as the proximal Llanos Basin.

Fault reactivation in Palaeocene time producedthree independent syn-orogenic basins separatedby low-amplitude uplifts with exposure of theCretaceous sedimentary cover that supplied sedi-ments to an axial depocentre. The Central Cordil-lera supplied sediments to the western and axialdomains, as recorded by detrital zircon populationages younger than 360 Ma, metamorphic rock frag-ments and garnet–epidote–hornblende heavy min-eral assemblage. In late Palaeocene–early Eocenetime, depositional systems of the axial depocentrewere bordered by structures associated with theinversion of extensional faults, which correspondto the border of the Cretaceous extensional basin.Cretaceous rocks exposed at both margins byinverted structures supplied sediments to the axialdepocentre. Maastrichtian–early Eocene rates ofsedimentation recorded in sections to the west ofthe axial depocentre have values similar to therates recorded in the western depocentre.

In addition to the Cretaceous sedimentarycover, another source area that supplied sedimentsto the eastern depocentre was a currently buriedbasement-cored high in the southern Llanos Basin,

as indicated by the dominance of a zircon popula-tion older than 1.2 Ga (Fig. 6), the immature compo-sition of upper Palaeocene sandstones, and a recordof unstable titanite and hornblende fragments. Intra-plate magmatism reached areas as far as basement-cored uplifts in the Llanos Basin, as indicated bythe presence of Palaeocene volcanic zircons withsediments supplied from such intraplate uplifts.

This research is part of a multidisciplinary project fundedby Ecopetrol S.A-ICP (project ‘Cronologıa de la deforma-cion de cuencas subandinas), Hocol S. A, Maurel & Prom,and Corporacion Geologica ARES. Thanks to J. Ortiz,G. Nova, P. Montano, O. Molsalve, J. Roncancio andJ. C. Caicedo, all of whom participated in the field workfor this project; to D. Ochoa, M. Romero, A. Pardo,S. daSilva, M. Paez and M. Rueda for their palynologicalanalysis; and C. Bustamente, C. Ojeda and A. Torres fortheir provenance analysis. We also want to thankN. Sanchez, J. Rubiano, N. Moreno (ICP-Ecopetrol) andM. Parra for beneficial discussions, and Y. Villafanez forthe construction of Figures 6 and 7. Our gratitude alsogoes to M. Ibanez-Mejia (University of Arizona) whoshared with us the geochronological data on basementuplifts in the proximal Llanos Basin. A. Reyes-Harker(Ecopetrol), M. de Freitas and A. Fajardo (Hocol) whogave us the permission for the publication of the data pre-sented in this paper. The manuscript was improved by edi-torial comments of two anonymous reviewers and editorM. Nemcok.

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onset of fault reactivation in the eastern cordillera of ...cmontes/camilomontes...palaeogene time...

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